The O, Sr, Nd and Pb isotope geochemistry of MORB

Chemical Geology, 62 (1987) 157-176 Elsevier Science Publishers B.V., Amsterdam - - Printed in The Netherlands

157

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THE O, Sr, Nd AND Pb ISOTOPE GEOCHEMISTRY OF MORB E M I ITO 1'.1, W I L L I A M M. W H I T E 2'*~ and C H R I S T A G ( ) P E L "~ ~Department o[ Geology and Geophysics, University of Minnesota, Minneapolis, MN 55455, and Department of Terrestrial Magnetism, Carnegie Institution of Washington, Washington, DC 20015 ( U.S.A.) ~Max-Planck~Institut fi~r Chemie, D-6500 Mainz (Federal Republic o[ Germany) and College of Oceanography, Oregon State University, Corvallis, OR 97331 (U.S.A.) :~Laboratoire de Gdochimie et Cosmochimie, Institut de Physique du Globe, Universitd de Paris, F 75230 Paris C~dex 05 ( ~¥ance) (Received February 28, 1986; revised and accepted August 11, 1986)

Abstract Ito, E., White, W.M. and GSpel, Chr., 1987. The 0, Sr, Nd and Pb isotope geochemistry of MORB. Chem. Geol., 62: 157-176. We report analyses of 0, Sr, Nd and Pb isotope ratios in 52 fresh glasses of mid-ocean ridge basalt ( MORB ). JlsOvalues in basalts range from +5.35 to +6.05%o and do not correlate with indices of fractional crystallization or with STSrff6Sr. Values up to + 6.47%c occur in andesites and dacites from the propagating rift at 95°W on the Galdpagos Spreading Center. Neither fractional crystallization nor assimilation of altered material appear to be dominant controls on the ~ 1sO of basalts. A weak, but statistically significant correlation was found between J 180 and Nd and Pb isotope ratios in Pacific MORB which warrants further investigation. Otherwise, no statistically significant correlations between J ~ 0 and radiogenic isotope ratios were h)und, indicating that any 0 isotope variations in the depleted upper mantle are largely unrelated to incompatible-element enrichments and depletions. There are important and significant differences in variances of radiogenic isotopes in MORB from different ocean basins, but these differences do not appear to be related to spreading rate. We conclude that they reflect real differences in the homogeneity of the upper mantle beneath ocean ridges and that efficacy of mixing during magmatic processes is not a dominant control of the degree of heterogeneity observed in MORB. Pacific MORB has more uniform isotope ratios than MORB from other ocean basins. Radiogenic isotope ratios are most strongly correlated among Pacific MORB and least correlated among Indian Ocean MORB. Means of radiogenic isotope ratios of MORB from various ridges also vary considerably. However, there is no consistent relationship between means of various isotope ratios: means of both STSr/S~Sr and 14:~Ndp44Ndare lower in Pacific than in Atlantic MORB; Pb isotope ratios in MORB from the two oceans are similar. Indian Ocean MORB has higher STSr/S~Srand lower ~4:/Nd/'44Ndthan MORB from the Atlantic and Pacific but lower 2°~Pbff°4Pb. Our data strongly reinforce the view that the depleted upper mantle consists of a number of distinct subreservoirs produced by a variety of processes. Although the Pb presently in the upper mantle may not have long resided there, it could not have been derived from a primitive reservoir. A positive correlation between 2°6Pbff°4pb and 2°sPb*ff°6Pb* suggests the Pb in the mantle beneath the Pacific and Atlantic is derived from an enriched reservoir. The Pb isotope ratios in Indian Ocean MORB are suggestive of the presence of a recycled component.

Present addresses: *~Department of Geology and Geophysics, University of Minnesota, Minneapolis, MN 55455, U.S.A. *2Department of Geological Sciences, Cornell University, Ithaca, NY 14853-1504, U.S.A.

0009-2541/87/$03.50

© 1987 Elsevier Science Publishers B.V.

158

1 .Introduction Mid-ocean ridge basalt (MORB) is certainly the most abundant volcanic rock-type, and it seems reasonable to assume that the source region of MORB must represent a significant portion of the upper mantle. A thorough knowledge of the isotopic composition of MORB and their source would thus seem a prerequisite for understanding crust-mantle evolution, particularly in view of the abundant trace-element and isotopic evidence that the MORB source is the incompatible-element-depleted complement of the continental crust. Of the numerous isotopic studies which have been made of MORB, several have concentrated on apparently geochemically anomalous regions of ridges associated with oceanic islands (e.g., White and Schilling, 1978; Dupr~ and All~gre, 1980; Verma and Schilling, 1982 ) ; others have provided data suggesting a number of interesting features in the isotope geochemistry of MORB, including apparent correlations between various isotopic ratios and possible inter-oceanic differences in isotopic compositions (e.g., Cohen and O'Nions, 1982a; Dupr~ and All~gre, 1983). If, as is often assumed, the oxygen isotopic fractionation among minerals in the mantle is nearly zero, there is little a priori reason to assume that any O isotopic variation will exist in the mantle. However, studies by Garlick et al. (1971) and Kyser et al. (1982) have demonstrated the existence of such variations. The hypothesis of Kyser et al. suggests prior melt extraction should have resulted in differences in 51s0 between the depleted MORB source and enriched or undepleted plume sources, but the possibility that these O isotopic variations are related to variations in other isotopic and chemical parameters has not been adequately investigated. Another possible cause of source-related variations in O isotopic composition would be the presence of recycled continental or oceanic crustal material. The presence of such a component: has been suggested on the basis of the

radiogenic isotopic signatures of some oceanic island basalts (Cohen and O'Nions, 1982b; White and Hofmann, 1982; Dupr~ and All~gre, 1983). Continental material and sediments have greatly elevated 5~80; reservoirs in the mantle containing such material could be expected to have different ~lsO as well as different Sr, Nd and Pb isotope ratios. Recycling of altered oceanic crust (Chase, 1981; Hofmann and White, 1982) might also result in either upward or downward shifts of 3180 depending on whether alteration occurs primarily at low or at high temperature (e.g., Gregory and Taylor, 1981 ). We undertook an extensive study of O, Sr, Nd and Pb isotopic compositions in MORB with the following objectives in mind: (1) to provide a baseline data-set of N-type MORB (defined on p. 161 ) and determine how often basalts with radiogenic Pb and Sr and unradiogenic Nd occur on ridge segments far from hot spots; (2) to determine to what degree O isotope ratios vary in MORB, and whether such variations are related to Sr, Nd and Pb isotopic compositions or other chemical parameters; (3) to investigate whether very small amounts of hydrothermal alteration or weathering of MORB samples has any effect on the isotopic parameters; ( 4 ) to determine to what degree variations in different radiogenic isotope ratios are correlated; (5) to determine whether there are interoceanic differences in either means or variability of isotopic ratios; and (6) seek evidence of recycled oceanic crust or sediments in the O, Sr, Nd and Pb isotopic composition of MORB. Finally, we sought to examine relationships between isotopic and trace-element compositions of MORB. This aspect, and results of our trace-element studies, are discussed in Hofmann and White (1983), Jochum et al. (1983), and Hofmann et al. (in prep.).

2. S a m p l e s a n d a n a l y t i c a l m e t h o d s We selected 52 fresh, glassy samples for analysis; 17 from the Atlantic Ocean, 27 from the

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Pacific Ocean, and 9 from the Indian Ocean (Fig. 1 ). Most came from the collections of the Smithsonian Institution, and major-element composition of these samples is reported in Melson et al. (1977), or is available from National Museum of Natural History, Smithsonian Institution (T. O'Hearn). A number of samples from the Pacific ( sample designations beginning with K ) was provided by H. Puchelt. Major-element analyses for those samples are reported in Puchelt and Emmermann (1983). Trace-element chemistry of some samples is given in Puchelt and Emmermann (1983) ; data on other samples are or will be reported elsewhere (Hofmann and White, 1983; Jochum et al., 1983; Newsom et al., 1986; Hofmann et al., in prep. ). The glass chips were hand-picked after crushing to ~ - 50 mesh, ultrasonically cleaned with 7 N HNO3 solution and distilled water, and then dried, examined and selected using a binocular microscope for the second time. At this stage, any chips containing large phenocrysts were removed. An aliquot of the clean chips was ground using an agate or alumina mortar and pestle for 0, Sr and Nd isotopic and trace-element analysis, and the remainder was saved as chips for Pb isotopic analysis. For O isotopic analysis, the oxygen was extracted using BrF5 (Claytol~ and Mayeda, 1963 ). All the preparatory steps within the vacuum apparatus, prior to the actual reaction with

the reagent, were performed at ambient temperature as the glass powders proved to be highly reactive. The CO2 gas was analyzed with either a double-collector 60 ° sector, or a Nuclide ® triple-collector 90 ° sector, mass spectrometer at the Geophysical Laboratory in Washington, D.C. Sr and Nd isotopic analyses were performed on Finnigan ® MAT261 mass spectrometers at the Max-Planck-Institut ftir Chemie (M.P.I.) in Mainz. Analytical procedures for Sr and Nd have been described in White and Patchett (1984). Pb isotopic analysis of 14 of the samples was performed on a Cameca ® THN206 instrument in the Institut de Physique du Globe, University of Paris (I.P.G.), and followed the procedures of Manh~s et al. (1978). Generally, 0.5-1.0 g of sample was used and blanks were in the range of 500-700 pg. The remainder of the Pb analyses were performed at M.P.I. on a MAT261 mass spectrometer equipped with multiple Faraday cups for simultaneous collection of all Pb isotopes. Analytical procedures were adapted from Manh~s et al. (1978) and are described in White and Dupr~ (1986). During the course of this study (1981-1985), blanks decreased from 220 to 150 pg for Sr, 170 to 60 pg for Nd, and 400 to 100 pg for Pb. Analytical errors based on repeated measurements of standards are given in the footnote to Table I; internal errors were virtually always better. Most of these samples were analyzed more than once (only means are reported here), hence the actual analytical uncertainty is generally better than the errors given in Table I, which pertain to single analysis. 3. R e s u l t s

Results are listed in Table I. Our intention in this study was to focus on samples whose chemistry is not affected by mantle plumes or hot spots; such samples will hereafter be termed normal- or N-MORB (following Schilling, 1975). We distinguish them from Schilling's other classes (T- and P-MORB) solely on a

160 TABLE ] N e w Sr, N d , P b a n d 0 isotope d a t a on M O R B Location

~TSrff~Sr

14:~Nd/144Nd

'-'('6Pbff°aPb

2°TPbff°4pb

-"'SPb/'-'"'Pb

(~lt~O (!:&,}

+5.55 +5.54 +5.55 +5.60 +5.97 +5.61 ~5.55 +5.35 +5.80 +5.70 +5.94 +5.67 +5.75 +5.90 +5.86 + 5.60 +5.69 +5.62 +5.61 +5.69

ATLANTIC OCEAN

VG962 VG367 VG965 VG200 521-1B 528 3 534 2 1 VG968 VG744 VG205 VG296 VG937 VG94I VG249 VG260 P6906 28B GS7309-94 GS7309-75 VG198 VG192

70.17°N 52.67=N 49.81 ° N 42.96°N 36.82°N 36.81 " N 36.80°N 28.90°N 25.40~N 22.92°N 22.24 N 22.24°N 22.17°N 11.22~N 11.02~N 6.01 ° N 0.02°S 0.55°S 21.87°S 21.93-$

15.26°W 34.94°W 28.65°W 29.20°W 33.27°W 33.26°W 33.27°W 43.32°W 45.30°W 13.51°W 45.02°W 45.02°W 45.25°W 43.06°W 43.67°W 33.28 ° W 24.58°W 16.07°W 11.85°W ll.81°W

0.70315 0.70316 0.70285 0.70334 0.70290 0.70285 0.70288 0.70287 0.70261 0.70281 0.70232 0.70250 0.70237 0.70245 0.70253 0.70276 0.70261 0.70255 0.70229 0.70232

0.513169 0.513146 0.513209 0.512992 0.513077 0.513066 0.513079 0.513062 0.513145 0.513130 0.513214 0.513220 0.513185 0.513149 0.513127 0.513050 0.513115 0.513100 0.513182 0.513182

18.343 18.344 18.339 19.690 18.814 18.846 18.899 18.589 18.320 18.275 18.317 18.338

15.503 15.475 15.499 15.608 15.541 15.534 15.545 15.529 15.501 15.485 15.485 15.481

37.792 37.870 37.830 39.299 38.404 38.361 38.435 38.108 37.807 37.842 37.710 37.700

18.408 18.359 19.444 18.845 18.775 18.375 18.299

15.490 15.504 15.588 15.575 15.568 15.518 15.489

37.906 37.845 39.037 38.256 38.320 37.842 37.726

46.39°N 44.66°N 44.66° N 44.66°N 44.66°N 44.27°N

130.22°W 130.33°W 130.33°W 130.33°W 130.33°W 129.08°W

0.70244 0.70258 0.70256 0.70256 0.70249 0.70249

0.513030 0.513069 0.513176 0.513145 0.513165 0.513250

18.766 18.751 18.486 18.448 18.471 18.470

15.547 15.563 15.485 15.466 15.500 15.457

38.226 38.568 37.852 37.786 37.903 37.867

+5.81 +6.04 +5.64 +5.79

13.83°N 12.14~N 1.45°N 3.78°S 13.22°S 20.36°S 31.00°S 20.39°S

104.14°W 103.83°W 101.35°W 102.73°W 112.33°W 114.02°W 113.12°W 113.78°W

0.70256 0.70248 0.70249 0.70244 0.70242 0.70246 0.70252 0.70245

0.513130 0.513121 0.513182 0.513193 0.513180 0.513161 0.513097 0.513156

18.336 18.337

15.495 15.501

37.837 37.840

+5.62 +5.49 +5.57

17.975

15.442

37.478

18.321 18.616 18.408

15.484 15.518 15.490

37.798 37.957 37.906

2.18~N 2.08°N 2.62°N 2.70°N 2.70°N 2.70°N 2.70°N 2.70°N 0.74°N 0.71°N 1.04°N 1.04°N 1.45°N

100.67°W 100.34°W 95.28°W 95.24°W 95.24°W 95.24°W 95.24°W 95.24°W 85.58°W 85.50°W 85.12°W 85.12°W 85.10°W

0.70247 0.70243 0.70283 0.70280 0.70282 0.70283 0.70280 0.70281 0.70254 0.70313 0.70277 0.70249 0.70240

0.513124 0.513157 0.513040 0.513043 0.513062 0.513075 0.513046 0.513055 0.513101 0.513119 0.513118 0.513142 0.513198

18.269 18.228 18.744 18.743 18.741 18.749 18.736 18.752 18.574 18.554 18.644 18.366 18.287

15.479 15.453 15.562 15.540 15.559 15.560 15.550 15.564 15.515 15.558 15.548 15.500 15.481

37.797 37.697 38.566 38.548 38.552 38.550 38.516 38.568 38.132 38.215 38.236 37.911 37.816

PACIFIC OCEAN

Juan de Fuca: VG768 VG44 DIO 1 DlO-2 D10-3 VG348

+5.46

EPR: R3-3-D30 R3-3-DIO VG973 K28a-D232E VG875 KIO-D33A VG798 K12A-D33Aa

+5.56 +5.61 +5.89

GSC: K42a-D206F K46a-D202F K62a-D143G VG1235 VGI214 VG1234 VG1202 VG1223 K71a-D130H VGIO01 VG1770 VG1747 K73a-D123H

+5.71 +5.72 +5.99 +6.20 +6.17 +6.47 +6.12 +5.92 +5.42 +5.66 +5.80 +5.81

161 T A B L E I (continued) Location

S7Sr/SeSr

14:~Nd/144Nd

~°~Pb/2fJ4pb

2(~Tpbff°4pb

~'~Pbff°4pb

d 1sO ( %~

0.70295 0.70283 0.70274 0.70284 0.70276 0.70274 0.70303 0.70304 0.70311

0.513143 0.513074 0.513096 0.513114 0.513128 0.513109 0.513072 0.513083 0.513077

18.084 17.978 18.009 17.997 18.100 18.170 17.315 17.325 17.469

15.452 15.451 15.473 15.460 15.474 15.500 15.443 15.456 15.449

37.804 37.760 37.846 37.816 37.900 38.064 37.251 37.287 37.456

+5.72 +5.48 +5.58 +5.70 +5.74 +5.80 +5.97 ~6.03

INDIAN OCEAN

VG1583 VG5262 VG5269 VG5284 VG5294 VG5291 VG3095 Al193-I1-I03 Al193-15-23

5.35°N 3.78°N 3.70°N 1.65°S 5.28~S 5.36°S 24.98°S 24.98°S 25.78°S

68.69°E 63.87°E 63.89°E 67.77°E 68.53°E 68.62°E 69.99°E 70.01°E 70.23°E

SVSr/S~Sr ratios are normalized to 8 6 / 8 8 = 0 . 1 1 9 4 0 and are relative to SVSrff~Sr-0.70800 for the E i m e r & A m e n d ~ standard. E s t i m a t e d analytical errors are _+0.000035 ( 2 a errors) based on replication of s t a n d a r d s (in-run errors were generally b e t t e r ) . 1'~*Nd/t44Nd ratios are normalized to 146/144=0.72190 and are relative to values of ~taNd/~44Nd-0.511847 and 0.512647 for the La Jolla and B C R - I s t a n d a r d s respectively. E s t i m a t e d analytical errors are _+0.000020. Pb isotope ratios were corrected for fractionation using 1.40!i( per a.m.u, correction and are relative to values of 2°6PbF°4pb = 16.937, 2°7pb/2°4pb = 15.493 a n d 2°~Pb/2°4pb = 36.705 for the NBS-98I standard. Analytical uncert a i n t i e s (2a) are e s t i m a t e d at ~°6Pbff°4pb= _+0.010, 2°TPbF°4Pb=_+0.016 and 2°SPbff°4pb= _+0.035. 0 isotope ratios were normalized against S L A P ( 6 ~80 = - 55.0%~ ). The 6 ~sO for NBS-28 Quartz on this scale using our e q u i p m e n t and procedures is ~ 9.50 _+.0.10~" ....

geographic basis. All samples from known hotspot- or plume-affected ridge segments and their associated geochemical transition zones (e.g., Hart et al., 1973; White et al., 1976; Verma and Schilling, 1982) are considered to be P-type, regardless of composition. All samples, regardless of composition, not from ridge segments influenced by known plumes or hot spots, are considered by us to be N-MORB. In other words, samples from ridge segments between hot spots and the point where the geochemical signal of the hot spot decays to the baseline noise level are P-type; all others are not. By this criterion, only VG200 from 43 ° N, the samples from 95 ° W on the Galfipagos Spreading Center (GSC), and the Three FAMOUSsamples (5211B, 528-3, 534-2-1) are P-type (in Schilling's original classification, the GSC and FAMOUS samples are T-type). We note, however, that the 95 °W Galfipagos Spreading Center ( GSC ) samples are at the limit of the influence of the Galfipagos plume and our classification of them as P-type is somewhat arbitrary. Fig. 2 shows the variation of 87Sr/S6Sr, 2°6pb/2°4pb and 8180 along the Mid-Atlantic Ridge (MAR). Our data support the view that high 87Sr/S6Sr and 2°6pb/z°4pb ratios are generally restricted to the Iceland, 45°N, Azores

and 35 ° N anomalous regions identified by Hart et al. (1973), White et al. (1976), and Dupr6 and All~gre (1980). Nevertheless, there is considerable variability in isotopic composition outside these areas and 2°~Pbff°4pb ratios as high as 19.4 occur. This is also true of the Pacific (Verma and Schilling, 1982; Eaby et al., 1984; Macdougall and Lugmair, 1986; White et al., 1987) and the Indian Ocean Ridges (Hamelin and All~gre, 1985; Hamelin et al., 1986). However, among N-MORB, the Pacific ridges are more uniform than the Atlantic, a point which we shall subsequently discuss in greater detail. ~lSO-values in our samples range from +5.35 to +6.47%c, but most values are in the narrower range of +5.55 to +5.90%~. Fig. 2 indicates there is no systematic variation in (~180 along the Mid-Atlantic Ridge comparable to that seen in radiogenic isotopes. In particular, VG200 from the 45 ° N area, which has the highest 87Sr/S6Sr and 2°6pbff°4Pb and lowest 14:~Nd/144Nd in our data set, does not have unusual 6 ~sO. d~SO-values are plotted against Mg-number [ = Mg/( Feto t + Mg) ] and SiO2 in Fig. 3. Of the samples with ~lsO higher than + 5.90%~, four are andesites and dacites from the propagating rift at 95°W on the GSC. In these cases, the

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Fig. 2. Variation in 5~sO, STSr/S~Sr and 2°~Pbff°4Pb along the Mid-Atlantic Ridge. Solid symbols=our data; open symbols=data from Hart et al. (1973), Sun et al. (1975), Hart (1976), O'Nions et al. (1977), Tatsumoto (1978), White and Schilling (1978), Cohen et al. (1980), Dupr6 and All~gre (1980), Sun (1980), Machado et al. (1982), White and Hofmann (1982), and Hamelin et al. (1984). Also shown is the range in isotope ratios in the FAMOUS region (data from White and Bryan, 1977; Kyser, 1980; Dupr~ et al., 1981 ). There are no systematic variations in (~180 corresponding to those seen in Sr and Pb isotope ratios.

51sO enrichment is logically attributed to the extensive fractional crystallization these sampies have experienced (Muehlenbachs and Byerly, 1982). Of the remainder with 51SO-values of above + 5.90%o, only K71a-D130H, from 85 °W on the GSC, has a major-element composition suggesting it has experienced more than usual fractional crystallization (Mg-number -- 44; Puchelt and Emmermann, 1983 ). Fig. 3 shows there is no systematic variation of 51so with Mg-number or SiO2 among the basalts. Fe-Ti-oxides do have significantly lower 51so than the liquids from which they crystallize ( Anderson et al., 1971; Taylor, 1968) ; the onset of oxide fractional crystallization results in a rise in both SiO2 and 51so.

Fig. 3. Variation of 5]sO vs. Mg-number and SiO~ (majorelement data from Melson et al., 1977; Puchelt and Emmermann, 1983; T. O'Hearn, unpublished data, 1983). Symbols: circles=Atlantic Ocean; squares=Pacific Ocean; triangles = Indian Ocean.

4. C o r r e l a t i o n s a m o n g isotope ratios One of the principal objectives of this study was to seek possible correlations between 5zso and radiogenic isotope ratios which might provide clues to the origin of chemical heterogeneity in the mantle. The question of the degree to which various radiogenic isotope ratios correlate in MORB has been controversial. We (White et al., 1981) have argued that there is no correlation between Sr and Nd isotopes among N-MORB; Cohen and O'Nions (1982a) and Macdougall and Lugmair (1986) have argued there is little correlation between these ratios among Pacific MORB; Hamelin et al. (1984) argue 2°6pb/2°4pb and STSr/SGSr correlate in both Pacific and Atlantic MORB, but they (Hamelin et al., 1986) note little correlation exists among Indian Ocean MORB. In this section, we examine these questions in a more quantitative manner using all available highquality data. Figs 4-7 are plots of various isotope ratios against one another. Several features of these diagrams merit brief comment. Fig. 4 confirms the tendency, noted by Hamelin

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P-o6pb/2O4pb Fig. 4. Variation of 2°Tpbff°4pb and 2°SPbF°4pb with 2°~Pbff°4pb in MORB. Symbols: circles=Atlantic Ocean; squares=Pacific Ocean; triangles=Indian Ocean; solid symbols are our data, open symbols are from references given in Fig. 2 and Church and Tatsumoto (1975), Duprd and Allbgre (1983), Hamelin and Allbgre (1985), Hamelin et al. (1986), Price et al. (1986) and White et al. (1986). Note that Indian Ocean MORB plot to higher 2°Tpb/2°4pb and 2°SPbff°4Pb than Atlantic and Pacific MORB.

and All~gre (1985) and Hamelin et al. (1986), of Indian Ocean MORB to plot at higher 2°Spbff°4pb and 2°TPbff°tPb than Pacific and Atlantic MORB. Our data suggest this is a feature of all Indian Ocean MORB, not just those of the Southwest (SWIR) and Southeast (SEIR) Indian Ocean Ridges. For the most part, the Indian Ocean 2°SPbff°tpb-2°6pbff°4pb data form a strongly correlated linear trend with shallower slope than the Pacific and Atlantic data; and the Indian Ocean e°Tpbff°4pb2°6Pbff°4pb data show considerably more scatter than the Atlantic and Pacific data. Fig. 5 shows STSr/S6Sr and 143Nd/lt4Nd are well correlated when all data are considered. The correlation is poorer when only N-MORB are considered, but is nonetheless significant, as we

Fig. 5. Variation of 14aNd/144Nd and S~Sr/S~Sr in MORB. Symbols are as in Fig. 4. The correlation between these ratios is poorer in Indian Ocean MORB than in Atlantic and Pacific MORB. Atlantic MORB tend to plot above (to higher 143Nd/144Nd and STSr/S~Sr) the Pacific data. Sources of data (in addition to Table I and references given in caption to Fig. 2) are LeRoex et al. (1983), Verma (1983), Verma et al. (1983), Eaby et al. (1984), Hamelin and All~gre (1985), Hamelin et al. (1986), Macdougall and Lugmair (1986), Price et al. (1986), and White et al. (1986).

shall subsequently show. The Pacific and Atlantic data are better correlated than the Indian Ocean data. There is also a slight tendency for Atlantic MORB to have higher mNd/144Nd than Pacific MORB, a point upon which we shall elaborate in the following section. Fig. 6 shows STSr/S~Sr and 143Nd/ltnNdplotted against 2°6pbff°4pb. Overall, the correlations are weak or non-existent. This is largely due to the scatter in the Indian Ocean data. Indian Ocean MORB forms fields largely distinct from that of the Pacific and Atlantic on these plots. The correlation between 143Nd/144Ndand 2°~Pbff°4pb is better than that between 87Sr/S6Srand 2°~pbff°4pb. Fig. 7 shows STSr/S6Sr and 143Nd/144Nd plotted against 61sO. There appears to be little correlation between either ratio and oxygen isotopic composition. Table II gives the correlation matrices for NMORB for the three ocean basins. Whereas we

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have plotted all available data in Figs. 4-7, we give the correlation coefficients only for the NMORB subset in Table II, because inclusion of data from ridge segments influenced by hot spots or plumes may yield correlations which are the trivial result of mixing of hot-spot and N-MORB sources. Correlations which are statistically significant at the 1% level are shown in italicized typeface. A striking feature of Table II is the large disparity between oceans in the quality of the correlations. In the Pacific Ocean, all correlations among radiogenic isotope ratios are significant atthe ] % level; in the Indian Ocean only P b - P b correlations are significant. In the Atlantic data, correlations between 87Sr/S6Sr and Pb isotope ratios are not statistically significant at the 1% level, though the STSr/S6Sr-2°8pbff°4pb correlation is significant at the 2.5% level. In both the Atlantic and Pacific, correlations between 1 4 3 N d / 1 4 4 N d and Pb isotope ratios are

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better than between S7Sr/S6Sr and Pb isotope ratios. This is more apparent when the coefficients of determination are compared. This parameter, which is simply the square of the correlation coefficient, "determines" the proportion of variability of one variable which is related to the variability of the other variable. Thus, in the Atlantic and Pacific, only ~ 10% (0.3 squared) of the variation in S7Sr/S6Sr is related to variation in 2°6pbff°4pb, while ~ 50% (0.7 squared) of the variation in 143Nd/144Nd is related to variation in 2°6pbff°npb. Variations in 87Sr/S6Sr and 1 4 3 N d / 1 4 4 N d in the Indian Ocean mantle are virtually entirely independent of variations in other isotope ratios. As expected from Fig. 7, correlations between 51so and radiogenic isotope ratios are generally not statistically significant. The exceptions are the correlations with Nd and Pb isotope ratios among Pacific MORB, which are significant at

165 TABLE II Correlation matrices for N-MORB 14:~Nd/144Nd

2o6pbffO4pb

2oTpb/2O4pb

~oSpbffO4pb

- 0 . 3 4 3 (82)

0.346 (66) - 0 . 6 5 8 (55)

0.582 (66) - 0 . 7 5 3 (55) - O. 768 (91 )

0.573 -0.717 0.8 73 0.848

dlsO

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SVSr/S6Sr '4:~Nd/'44Nd -'°6pb/~°4pb ~°TPbff°4pb 2°spb/2°4pb

(66) (55)

-0.342 (23) - 0 . 6 8 3 (23)

(91 )

0.594 (21 ) 0.540 ( 21 ) O. 725 ( 21 )

(91)

Atlantic Ocean:

~:Sr/SSSr ~4"~Nd/144Nd 2°6pb/2°4pb 2°Tpbff°4pb '-'°SPbff°4pb

- 0 . 5 2 6 (34)

0.215 (38)

0.062 (38)

0.267 (38)

- 0 . 7 5 9 (20)

- 0 . 5 8 4 (20) 0.781 (60)

- 0 . 7 4 3 (20) 0.970 (60) 0.801 (60)

-0.247 (41) -0.095 (36)

0.109 (41) -0.111 (36) 0.545 (50)

-0.147 (41) -0.165 (36) 0.850 (50) 0.621 (50)

-0.451 0.240 -0.253 -0.292 -0.278

(19) (19) (16) (16) (16)

0.679 -0.163 -0.738 -0.112 -0.668

(8) (8) (8) (8) (8)

I n d i a n Ocean:

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-0.211 (36)

Number of samples shown in parentheses. Correlations which are statistically significant at the 1% level are shown in italicized type. Data filtered as described in text. See captions of Figs. 4-6 for data sources.

the 1% level. In the Pacific, there is a tendency for "enriched" radiogenic isotope signatures to be associated with high glSO-values (a significant positive correlation exists between g180 and SVSr/S6Sr if three samples are excluded). While these correlations are statistically significant, we view them with caution and skepticism for two reasons: (1) they depend rather heavily on a few extreme points; and (2) the correlations are not seen in the data from the Indian and Atlantic Oceans. Indeed, the Indian and Atlantic Ocean correlation coefficients for O and Pb isotope ratios have the opposite sign to those for the Pacific. However, we have already noted that correlations vary markedly between ocean basins, and the higher correlations of ~lso with radiogenic isotope ratios is consistent with the overall higher correlations

in the Pacific data between radiogenic isotope ratios. Thus we are not willing to entirely dismiss these correlations. Correlations between isotope ratios tend to be better when only data from a single ocean basin are considered than when the entire data set is considered. This appears to be an extension of the observations of Macdougall and Lugmair (1986) and White et al. (1987), who noted that correlations between STSr/S6Sr and 143Nd/144Nd and between Pb isotope ratios are generally better on regional scales (kilometers to hundreds of kilometers) than they are for Pacific MORB generally. Correlations between isotope ratios presumably reflect coherent behavior of parent-daughter pairs during fractionation events. Different processes may involve different behaviors of these par-

166 ent-daughter pairs. Where several events and processes have contributed to mantle heterogeneity, correlations among isotope ratios are likely to be poorer than where only one event or one process is involved. The general decline in the quality of isotopic correlations with increasing scale of sampling suggests that, locally, the mantle may be affected by a small number of events and/or processes, whereas globally, a larger number of events/processes are involved.

5. Inter-oceanic comparisons, isotopic diversity and spreading rates We have seen in the previous section that the quality of correlations among isotope ratios in N-MORB varies from ocean to ocean. Cohen and O'Nions (1982a), and Hamelin et al. (1984) have suggested isotopic compositions of MORB from the East Pacific Rise (EPR) are more homogeneous than those of the MAR. Batiza (1984), and Hamelin et al. (1984) have speculated these differences are related to spreading rate. Hamelin and All~gre (1985), and Hamelin et al. (1986) have noted the unique isotopic character ofbasalts from the SWIR and SEIR, including low 2°6Pbff°4pb and high 2°TPbff°tpb and 2°sPbff°tpb relative to 2°~Pbff°4pb. All~gre et al. (1984) argue that, in addition to the inverse relationship between isotopic diversity and spreading rate, mean values of STSr/S~Sr and Pb isotope ratios correlate positively with spreading rate. In this section, we examine possible inter-oceanic differences in variability and mean isotopic composition in the light of the large amount of data which has recently become available (e.g., Hamelin and All~gre, 1985; Hamelin et al., 1986; Macdougall and Lugmair, 1986; Price et al., 1986; White et al., 1987; this work). Table III gives means, standard deviations, and numbers of analyses of Sr, Nd, Pb and O isotope compositions for various ridges and ocean basins. We list the parameters for both the entire data set and a version (labelled N-

MORB) filtered to remove samles from plumeor hot-spot-influenced ridge segments. This is important as there is some bias in the data set toward plume-related segments as a result of intensive studies of such segments (e.g., White and Schilling, 1978; Dupr~ and All~gre, 1980; Verma and Schilling, 1982). The filtering criterion was discussed on p. 161. Table III confirms a number of inferences made in previous studies: MORB from the EPR, and indeed all Pacific ridges, are isotopically more homogeneous than those from the MAR and this is true regardless of whether one compares filtered or unfiltered versions of the data set. The fast spreading EPR (and indeed all Pacific ridges) has lower mean 87Sr/S6Sr than the MAR, again true of both the filtered and unfiltered data sets. Indian Ocean MORB has higher STSr/S6Sr and lower 143Nd/144Nd and 2°~Pbff°4Pb than the Pacific and Atlantic MORB. All these differences are statistically significant at or above the 5% confidence level and are broadly consistent with the postulates of Batiza ( 1984 ) and All~gre et al. ( 1984 ). There are no statistically significant differences in Table III does, however, reveal several surprises. First, the mean 2°6Pbff°4pb is identical (in a statistical sense) in N-MORB from the EPR and MAR. Most importantly, the mean 143Ndp44Ndratio of N-MORB from the fastspreading EPR is lower than the mean value for N-MORB from the slow-spreading MAR. The difference is statistically significant at the 0.5% confidence level and is opposite to that expected from the higher S7Sr/S6Sr in MAR basalts and from the general inverse correlation between 143Nd/144Nd and 87Sr/S6Sr. 143Nd/144Nd and STSr/S~Sr in Atlantic and Pacific MORB thus form two quasi-parallel arrays. This can be seen in Fig. 5, though the scatter in the data, particularly the Atlantic data, tends to obscure the difference. Another important feature of Table III is that the differences in 2°6Pbff°npb between Indian Ocean MORB on the one hand and Atlantic and

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Pacific MORB on the other are not accompanied by corresponding (statistically significant) differences in 2°!pb/2°4pb and 2°Spbff°4pb, confirming the conclusion of Hamelin and All~gre (1985) and Hamelin et al. (1986) that Indian Ocean MORB have higher 2°TPbff°4pb and 2°Spb/2°4pb for a given 2°6pbff°4Pb than Atlantic and Pacific MORB. It is important to note that this observation is true of the Carlsberg-Central Indian Ridge (CIR) as well as the SWIR and SEIR. 2°Tpb/2°4pb means and standard deviations are virtually identical for all ridges. This may in part reflect the relatively large contribution (perhaps one-quarter to one-third) of analytical error to the 2°Tpb/2°4pb variances, but also suggests isotopic differences between ocean basins have arisen largely in the latter part (the last ~ 2 Ga) of Earth history (because 2"~U, the parent of 2°TPb had largely died away by then). While there are significant differences among spreading centers in isotopic variability and mean isotope ratios, and while our compilation confirms the fast-spreading E P R show less isotopic variability than the slower spreading MAR, the compilation in Table III supports neither the hypothesis of an inverse relationship between isotopic diversity and spreading rate (Batiza, 1984), nor that of a relationship between mean isotopic composition and spreading rate (All~gre et al., 1984). This is further illustrated in Fig. 8, which is a plot of the standard deviation of STSr/S6Sr (SSr) against spreading rate. Following Batiza (1984), we have divided the E P R into three segments, the Rivera, Cocos and Nazca, which spread at very different rates. No significant correlation exists between spreading rate and Ss,, irrespective of whether one uses the filtered (N-MORB) or unfiltered versions of the data set. The relationships of the various Pacific and Indian Ocean ridges in Fig. 8 are particularly critical. There is, in effect, little variation in the standard deviation of STSr/S6Sramong all Pacific spreading centers, possibly excepting the GSC, which spreads at an intermediate rate and has

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a slightly higher standard deviation. The Gorda point could, perhaps, be justifiably dismissed on the basis of the very limited number of analyses (8), but there is no difference between the EPRNazca segment and the Juan de Fuca Ridge, between which spreading rate varies by nearly a factor of 3. There is presently sufficient data from the Juan de Fuca Ridge (Eaby et al., 1984) that the statistical parameters are unlikely to greatly change with additional data. Consider also the Indian Ocean Ridges. When all data are considered, the SWIR and SEIR have high variance, in part due to a few samples apparently related to the Crozet and New Amsterdam-St. Paul hot spots and having extremely radiogenic STSr/S6Sr (Hamelin and All~gre, 1985; Hamelin etal., 1986). When the samples from hot-spot-related segments are filtered out, the SWIR and CIR have similar standard deviations, both of which are lower than that of the SEIR, which spreads at a faster rate.

169

The slow-spreading CIR and SWIR are as homogeneous as the fast spreading E P R segments. These observations hold for ~43Nd/144Nd and 2°6Pbff°4pb as well as STSr/SGSr. There are still relatively few data from the Indian Ocean. Additional data may change the present picture; our point is only that the presently available data do not support the hypothesis of Batiza (1984). Batiza {1984) interpreted the apparent relationship between spreading rate and isotopic diversity as evidence that the mantle is everywhere heterogeneous, and that the degree of homogeneity observed in basalts depends on the efficacy of mixing during magmatic processes. He reasoned these mixing processes would be more effective at fast-spreading ridges than at slow-spreading ones, perhaps because of the presence of large magma chambers beneath the former and their absence beneath the latter (Sleep, 1975). Our compilation suggests real differences in the homogeneity of the upper mantle beneath ocean ridges, and that efficacy of mixing during magma processes is not a dominant factor controlling the degree of heterogeneity observed, though we agree that such processes may mask small-scale heterogeneity. Macdougall and Lugmair (1986) reached a similar conclusion. Our compilation suggests the mantle beneath all Pacific ridges is more homogeneous than that beneath the Atlantic, irrespective of spreading rate of the Pacific ridges. The upper mantle beneath the Indian Ocean also appears to be more homogeneous than the upper mantle beneath the Atlantic. The apparent greater homogeneity of the subPacific mantle is consistent with the more vigorous convective mixing beneath the Pacific suggested by All~gre et al. (1984). Further, mixing between depleted upper mantle and less depleted or enriched "blobs" derived from some deeper reservoir may well occur (indeed, such a process undoubtedly accounts for the more enriched isotopic signatures in MORB from ridges near hot spots). However, the lack of relationship between mean isotopic composi-

tion and the similarity of Pb isotope ratios in Pacific and Atlantic MORB and the lower 143Nd/144Nd is inconsistent with the model of depleted-mantle-enriched-blob mixing as formulated by All~gre et al. (1984). In summary, means and standard deviations of radiogenic isotope ratios of N-MORB from various ridges vary considerably on a global scale. However, there is no consistent relationship between the means of various isotope ratios: the Pacific contains the most "depleted" Sr but the most "enriched" Pb; the Atlantic contains the most "depleted" Nd; the Indian Ocean has the most "enriched" Sr and Nd but the most "depleted" Pb. Neither means nor standard deviations bear any systematic relationship to spreading rate. 6. C a u s e s of o x y g e n isotope v a r i a t i o n s Our data, as well as other data (e.g., Kyser, 1980), clearly demonstrate there are real variations in the oxygen isotope composition of MORB. An important question is the degree to which these variations reflect mantle source variations. All the following factors may affect the O isotope composition of MORB glasses: post-eruptional exchange with, or addition of, seawater; addition of seawater-altered basalt (low-temperature and/or hydrothermal) to the magma before and during eruption; 0 isotope fractionation during fractional crystallization and partial melting; and source heterogeneity. As stated earlier, all these glasses were handpicked and appear perfectly fresh. This precaution should rule out post-eruptional effects, though there is no guarantee that it does. Kyser and O'Neil (1984) concluded from hydrogen isotope compositions and water concentrations that addition of seawater to magma can occur and affects ~D-values. However, at the level of water addition they report (0.3%), this process will produce oxygen isotope changes of less than 0.05¢,~,. Assimilation of altered oceanic crust would produce correlated changes in STSr/S6Sr and ~lsO isotope ratios. For example, assimi-

170 lation of 10% hydrothermally altered crust having STSr/S~Sr of 0.705 and ~lsO of +3%0 would increase STSr/S6Sr by 0.00025 and decrease ~lsO by 0.25%0. The lack of correlation between STSr/S~Sr and ~lsO would argue against interaction with altered crust being an important influence on 0 and Sr isotope ratios, though small effects cannot be ruled out. Muehlenbachs and Kushiro (1974) and Kyser et al. (1982) have suggested that significant O isotope fractionation between minerals and melt occurs at magmatic temperatures. If so, both fractional crystallization and partial melting may lead to filSO-values of magmas which differ from those of their sources. This view has, however, been challenged (Gregory and Taylor, 1986). Our data argue against any large systematic effect of fractional crystallization on filsO until Fe-Ti-oxides crystallize and fractionate. We cannot rule out small effects, however, nor can we adequately assess the degree of possible fractionations during partial melting. The general lack of correlation between ~lsO and radiogenic isotope ratios indicates that if ~ s O variations in the MORB mantle source exist, they are largely unrelated to incompatible-element enrichments and depletions. On the other hand, there is a hint in the Pacific data of such a relationship. In our view, these correlations are too weak and based on too few data to attach great interpretive significance to them. Nevertheless, their statistical significance would appear to justify additional analytical work. If such correlations do exist, they may indicate that ancient magmatic events produced both variations in fi~sO (due to fractionations between minerals and melt) and parent/daughter ratios (which would in time lead to variations in radiogenic isotope ratios). Alternatively, mixing between depleted mantle and some other reservoir with different 3180 such as recycled material ( Hofmann and White, 1982) would lead to such correlations. We cannot be sure of exactly what causes the observed variations in O isotope ratios. None of

the factors we listed at the start of the section would appear capable of producing all the variations observed. We suspect that many, and perhaps all, of these contribute to the scatter. 7. C h e m i c a l e v o l u t i o n a n d f l u x e s i n t o t h e upper mantle

The evidence that MORB are derived from a reservoir depleted in the most incompatible elements is abundant and long-standing. Tatsumoto et al. (1965) noted the uniformly low incompatible-element (K, Rb, Th, U, Sr) abundances and 878r/S6Srratios in MORB but more importantly, pointed out Rb/Sr ratios were too low to generate the 87Sr/S6Sr observed. They concluded the MORB source had suffered a previous melt extraction which was related to continental crust formation. Since then, evidence from rare earths (e.g., Frey et al., 1968; Schilling, 1971 ), Nd isotopes (e.g., DePaolo and Wasserburg, 1976), and Hf isotopes (e.g., Patchett and Tatsumoto, 1980) have supported this conclusion. Nevertheless, it is clear the chemical evolution of the mantle has been complex. Isotope systematics, particularly Pb isotope ratios, of the mantle are not consistent with simple removal of melt to form the crust. Isotope ratios in oceanic basalts require a number of chemically distinct reservoirs in the mantle (e.g., White, 1985; Zindler and Hart, 1986). Heterogeneity of MORB sources is often ascribed to mixing with material derived from these other reservoirs (e.g., Schilling, 1973), or the inherent presence of such material in MORB source regions (e.g., Zindler et al., 1984). Others have suggested interaction with the core (e.g., Vidal and Dosso, 1978; All~gre et al., 1982) or recycling of material from the continental or oceanic crust as a cause of heterogeneity (e.g., Cohen and O'Nions, 1982b; Hofmann and White, 1982). Consideration of parent/daughter ratios and other trace-element ratios now indicates recycling of continental material into the regions of

171

the mantle sampled by oceanic volcanism must be very limited, particularly if this occurs by subduction of marine sediment (e.g., Patchett et al., 1984; White et al., 1984; Newsom et al., 1986). In particular, a consideration of U, Th and Pb abundances in marine sediments indicates ancient recycled sediment, when compared to the MORB mean, should now have low 2°6pb/z°4pb and high 2°7pb/2°4pb (White et al., 1984) and, hence, mixing of recycled sediment into the mantle cannot account for the distribution of Pb isotope ratios observed in MORB (e.g., Fig. 4) and, indeed, in most oceanic basalts (e.g., White, 1985; Fig. 1). Similarly, consideration of chalcophile- and siderophile-element abundances now indicates significant transfer of material from mantle to core is unlikely ( Newsom et al., 1986). In an important recent paper, Galer and O'Nions (1985) pointed out that T h / U ratios in the depleted mantle are much too low to account for the 2°8pbff°4pb-2°6pb/2°4pb systematics of MORB. The latter indicate Pb now in the depleted upper mantle must have evolved in a reservoir with a T h / U ratio (3.7-3.8) close to the bulk-Earth value, whereas Th isotope systematics indicate the depleted mantle presently has T h / U 2.5. Their interpretation is one of a steady-state upper mantle into which Th, U and Pb are being transferred from the lower mantle at the same rate these elements are being lost from the depleted upper mantle to the crust. They estimate a residence time for Pb in the upper mantle of the order of 800 Ma. In their model, Galer and O'Nions (1985) assume a bulk-Earth T h / U of 3.9. The case for such a value is not yet conclusive and others (e.g., Dupr~ et al., 1984 ) have argued for a T h / U as high as 4.2. A higher value would imply a longer residence time of Pb in the upper mantle. However, accepting for the moment the 3.9 value of Galer and O'Nions, we will briefly examine Pb isotope systematics of the upper mantle in light of their model. The model should apply to other decay systems such as S m - N d as well as to T h - U - P b . Fig. 9a shows

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2°Spb*ff°6Pb* (the ratio of radiogenic 2°Spb to radiogenic 2°6pb) plotted against 143Nd/144Nd. The weak but significant correlation could be interpreted as mixing between a highT h / U - l o w - S m / N d reservoir (such as primitive lower mantle), and a low-Th/U-high-Sm/Nd reservoir (depleted upper mantle), which is the relationship predicted by the model. Another and perhaps more fundamental prediction of the Galer and O'Nions model is an association of primitive 2°8pb*ff°6pb* ratios with primitive e°6pbff°4pb ratios. Using their values for primitive mantle, T h / U and 2asU/2°4pb of 3.9 and 8, respectively, the present primitive mantle 2°8pb*ff°6pb* and 2°6pbff°tpb should be ~0.96 and ~17.5, respectively. Fig. 9b shows 2°Spb~'ff°6Pb~ plotted against e°6pbff°4pb. The Galer and O'Nions (1985) model predicts a correlation between the two, one end of which should be defined by the

172

primitive mantle point. Overall, there is little correlation. The Pacific and the Atlantic data considered alone show a weak, but statistically significant correlation indicating an association of high 2°SPb~/z°6pb~ ratios with high 2°~pbff°4pb ratios. A comparison with Fig. 4 reveals high 2°6pb/2°tpb ratios belong to those points most displaced from the geochron and are therefore least primitive. Hence the observed association does not support the idea of a flux of Pb into the MORB reservoir from a primitive reservoir. The Indian Ocean MORB points scatter about the assumed primitive mantle point with many compositions plotting in the high-2°Spb~/Z°6Pb*,low-2°~pb/z°4Pb quadrant. The Indian Ocean MORB apparently evolved in a reservoir enriched in Th telative to U and depleted in U relative to Pb such as might be found in a reservoir containing ancient recycled sediment, but not in a primitive mantle reservoir. The conclusion we reach is that the Pb now in the depleted mantle cannot have been derived from a primitive reservoir, although we do not dispute the conclusion of Galer and O'Nions (1985) that the Pb now in the depleted mantle could not have long resided there. It might be argued that the reservoir from which Pb is derived is indeed primitive except for the postaccretional loss of Pb to the core, but the lack of correlation of abundances of siderophile and chalcophile elements, such as Mo and W, with Pb isotope ratios argues strongly against this hypothesis, because these elements would have been transferred to the core much more efficiently than Pb (Newsom et al., 1986). What is needed to explain the data is a reservoir with high Th/U, high U/Pb, high Rb/Sr and low Sm/Nd. Of the other reservoirs which might supply Pb to the depleted mantle, Galer and O'Nions (1985) have dismissed the mantle portion of the continental lithosphere on the grounds of its apparent stability and insufficient mass, and we, and they, have noted the inadequacies of recycling of continental material. A reservoir containing recycled oceanic

crust can also be dismissed because: (1) the enrichment in Th over U in the oceanic crust compared to the mantle appears minimal; and (2) hydrothermal processes will enrich the oceanic crust in U (Edmond et al., 1979) and perhaps deplete it in Th (Chen and Wasserburg, 1985). These observations combined with those of Galer and O'Nions (1985) appear to point to the existence of an unidentified enriched reservoir of substantial volume somewhere within the Earth, from which Pb is transferred into the upper mantle. Until further constraints are developed, a discussion of the nature of this reservoir would involve only speculation and be of little value. Although recycling of continental material, and particularly of marine sediment, cannot be the principal process producing chemical heterogeneity, it may account for the origin of some specific reservoirs. The low ~°6Pbff°4pb and 143Nd/144Ndand high 2°TPbff°4pb, 2°Spbff°4pb, and STSr/S6Sr of Indian Ocean MORB are precisely the isotopic signature expected of ancient subducted sediment. Hence we agree with Hamelin and All~gre (1985) that sediment recycling is a viable interpretation in this instance. We note the high 61SO-values ( ~ 6.0) in the two samples with the lowest 2°6Pbff°4pb (VG3095 and AII93-11-103) are consistent with this interpretation. Recycled sediment should also have distinct trace-element fingerprints. Trace-element studies of Indian Ocean MORB would provide an excellent test of this hypothesis.

8. S u m m a r y and conclusions (1) Fresh mid-ocean ridge basalt glasses have filSO-values in the range +5.35 to +6.05~i,~, with most samples falling in the narrower range of + 5.5 to + 5.9 %c. Values above + 6.1%~ occur only in andesites and dacites from propagating rifts. (2) The d lSO-values in basalts are not systematically related to indices of fractional crystallization, and this along with a lack of

173 correlation with 87Sr/S6Sr indicates fractional crystallization and assimilation of altered material are not dominant influences on the 6180 of MORB, at least until the onset of oxide fractionation. We cannot, however, rule out such processes producing small effects and the possibility that partial melting produces fractionation effects. ( 3 ) In general, there is no correlation between radiogenic isotope ratios and 6180. Statistically significant correlations between 6180 and Nd and Pb do occur in Pacific MORB. These correlations are too weak and based on too few data to provide a basis for interpretation, but they do warrant further investigation. No differences in mean 6180 have been found between MORB of different ocean basins. (4) Important differences exist in means and variability of radiogenic isotope ratios in MORB from different ocean basins. We confirm that Pacific MORB have lower 87Sr/S6Sr and are more homogeneous than Atlantic MORB. However, Pacific MORB have lower mean 143Nd/lntNd than Atlantic MORB. There is little difference in mean values of Pb isotope ratios in N-MORB of the two oceans. This appears to rule out a greater proportion of enriched material in the sub-Atlantic mantle. Indian Ocean MORB are distinct in all radiogenic isotope systems from Atlantic and Pacific MORB. Significant differences in the degree to which various isotope ratios correlate also exist between ocean basins. Among Pacific N-MORB, statistically significant correlations exist between all isotope ratios. At the other extreme, significant correlations exist only between Pb isotope ratios in Indian Ocean N-MORB. In general, 143Nd/144Nd correlates better with Pb isotope ratios than does STSr/S~Sr. (5) A relationship between isotope diversity and spreading rate is not supported by the data compiled in this study. (6) If Galer and O'Nions (1985) are correct in asserting that Pb now in the depleted upper mantle could not have long resided there, the relationship between 2°8pb*ff°~Pb* and

2°6pbff°4pb demonstrates that this Pb could not be derived from a primitive reservoir. The evidence would instead seem to point to the derivation of this Pb from an as yet unidentified enriched mantle reservoir. (7) Above all, our data suggest that the depleted upper mantle consists of a number of distinct sub-reservoirs which may be related to convection patterns. Dup% and All~gre (1983), Hart (1984) and Hamelin and Allggre (1985) have made similar points. Furthermore, our results and those in the literature indicate a variety of processes have operated upon the MORB source reservoir, or, alternatively but equivalently, the MORB source reservoir has mixed with reservoirs produced by a variety of processes. This point is similar to that made by White (1985). Mantle evolution cannot be considered strictly in terms of depletion and mixing with an hypothesized primitive reservoir. We emphasize that clear evidence for the existence of a primitive reservoir from isotope systematics is wholly lacking. He isotope data require only that some reservoirs be less depleted ( or, more properly, less degassed) than others. (8) Among MORB, sedimentary recycling can be tentatively associated only with Indian Ocean MORB. The nature of the processes producing heterogeneity in other reservoirs remains enigmatic.

Acknowledgements We thank W. Melson and the Smithsonian Institution for providing the bulk of the samples, and N.O.A.A. Marine Minerals Project (P. Rona, Director), H. Puchelt, R. Altherr, R. Batiza and J. Honnorez for additional samples. H. Voshage and A. Hofmann collaborated with us in the early phases of this work. We also thank T.C. Hoering and D. Rumble III for making the stable-isotope mass spectrometers at the Geophysical Laboratory available to us. J. Macdougall, F. Frey, J. Mahoney, J. Natland and B. Hamelin provided preprints and allowed

174

us to use unpublished data in our compilations. We thank H. Feldmann and S. Kielinczuk for technical assistance and J. Poppleton and M.J. Armburst for word processing. This work was supported by N.S.F. Grant OCE-8201077 and the Max-Planck-Gesellschaft zur FSrderung der Wissenschaft. The manuscript benefited from critical reviews by B. Hamelin, H. Taylor and S. Wickham. References All~gre, C.J., Duprg, B. and B%vart, 0., 1982. Chemical aspects of the formation of the core. Philos. Trans. R. Soc. London, Ser. A, 306: 49-59. All~gre, C.J., Hamelin, B. and Dupr6, B., 1984. Statistical analysis of isotopic ratios in MORB: the mantle blob cluster model and the convective regime of the mantle. Earth Planet. Sci. Lett., 71: 71-84. Anderson, Jr., A.T., Clayton, R.N. and Mayeda, T.K., 1971. Oxygen isotope geothermometry of mafic igneous rocks. J. Geol., 79: 715-729. Batiza, R., 1984. Inverse relationship between Sr isotope diversity and rate of oceanic volcanism has implications for mantle heterogeneity. Nature (London), 309: 440-441. Carlson, R.W., Macdougall, J.D. and Lugmair, G.W., 1978. Differential Sm/Nd evolution in oceanic basalts. Geophys. Res. Lett., 5: 229-232. Chase, C.G., 1981. Ocean island Pb: two stage histories and mantle evolution. Earth Planet. Sci. Lett., 52: 277-284. Chen, J.H. and Wasserburg, G.J., 1985. U, Th and Pb isotopes in hydrothermal fluids from the Juan de Fuca Ridge. EOS (Trans. Am. Geophys. Union), 66:929 (abstract). Church, S.E. and Tatsumoto, M., 1975. Lead isotope relations in ocean ridge basalts from the Juan de Fuca-Gorda Ridge area, N.E. Pacific Ocean. Contrib. Mineral. Petrol., 53: 253-279. Clayton, R.N. and Mayeda, T.K., 1963. The use of bromine pentafluoride in the extraction of oxygen from oxides and silicates for isotopic analysis. Geochim. Cosmochim. Acta, 27: 43-52. Cohen, R.S. and O'Nions, R.K., 1982a. The lead, neodymium and strontium isotopic structure of ocean ridge basalts. J. Petrol., 23: 299-324. Cohen, R.S. and O'Nions, R.K., 1982b. Identification of recycled continent-derived material in the mantle from Pb, Nd, and Sr isotope studies. Earth Planet. Sci. Lett., 61: 73-84. Cohen, R.S., Evensen, N.M., Hamilton, P.J. and O'Nions, R.K., 1980. U-Pb, Sm-Nd and Rb-Sr systematics of mid-ocean ridge basalt glasses. Nature (London), 283: 149-153.

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